Phosphorite hosted zinc and lead in the Sekarna deposit (central Tunisia)

ARTICLE

Phosphorite-hosted zinc and lead mineralizationin the Sekarna deposit (Central Tunisia)

Hechmi Garnit & Salah Bouhlel & Donatella Barca & Craig A. Johnson & Chaker Chtara

Received: 16 February 2010 /Accepted: 2 December 2011 /Published online: 17 December 2011# Springer-Verlag 2011

Abstract The Sekarna Zn–Pb deposit is located in CentralTunisia at the northeastern edge of the Cenozoic Rohiagraben. Mineralization comprises two major ore types: (1)disseminated Zn–Pb sulfides that occur as lenses in sedi-mentary phosphorite layers and (2) cavity-filling zinc oxides(calamine-type ores) that crosscut Late Cretaceous and EarlyEocene limestone. We studied Zn sulfide mineralization in theSaint Pierre ore body, which is hosted in a 5-m-thick sedimen-tary phosphorite unit of Early Eocene age. The sulfide miner-alization occurs as replacements of carbonate cement inphosphorite. The ores comprise stratiform lenses rich in sphal-erite with minor galena, Fe sulfides, and earlier diageneticbarite. Laser ablation–inductively coupled plasma mass spec-trometry analyses of sphalerite and galena show a wide range

of minor element contents with significant enrichment ofcadmium in both sphalerite (6,000–20,000 ppm) and galena(12–189 ppm). The minor element enrichments likely reflectthe influence of the immediate organic-rich host rocks. Fluidinclusions in sphalerite give homogenization temperatures of80–130°C. The final ice melting temperatures range from−22°C to −11°C, which correspond to salinities of 15–24 wt.% NaCl eq. and suggest a basinal brine origin for the fluids.Sulfur isotope analyses show uniformly negative values forsphalerite (−11.2‰ to −9.3‰) and galena (−16‰ to−12.3‰). The δ34S of barite, which averages 25.1‰, is 4‰higher than the value for Eocene seawater sulfate. The sulfurisotopic compositions are inferred to reflect sulfur derivationthrough bacterial reduction of contemporaneous seawater sul-fate, possibly in restricted basins where organic matter wasabundant. The Pb isotopes suggest an upper crustal leadsource.

Keywords Zn–Pb deposits . Sedimentary phosphorites .

Sekarna . Central Tunisia

Introduction

The major Tunisian Pb–Zn–(Ba–F) deposits, which occur incarbonate rocks of Mesozoic to Tertiary age, are thought tohave formed from orogenically driven brines that circulatedthrough crystalline rocks and then ascended and reactedwith fluids in overlying rocks during collision of the Euro-pean and African plates (Rouvier et al. 1985; Bouhlel 2005).Although the major deposits are found in calcareous rocks,evidence from other Pb–Zn deposits in the Diapir Zone,including the well-studied Bougrine deposit, suggests thatsalt diapirism played an important role in the emplacement

Editorial handling: B. Lehmann

H. Garnit (*) : S. BouhlelDepartment of Geology, Faculty of Sciences of Tunis,El Manar University,2092 Tunis, Tunisiae-mail: [email protected]

S. Bouhlele-mail: [email protected]

D. BarcaDepartment of Earth Sciences, University of Calabria,via Ponte Bucci 4, Cubo 15B,87036 Arcavacata di Rende (CS), Italye-mail: [email protected]

C. A. JohnsonU.S. Geological Survey,MS 963, Box 25046, Denver, CO 80225, USA

C. ChtaraGroupe Chimique Tunisien (G.C.T.),110, Rue Habib Chagra,6002 Gabes, Tunisia

Miner Deposita (2012) 47:545–562DOI 10.1007/s00126-011-0395-y

of the deposits (Orgeval et al. 1989; Sheppard et al. 1996;Bouhlel et al. 2009).

The Sekarna Zn–Pb deposit is distinctive in that it occurs insedimentary phosphorites of Early Eocene age. The settingthus differs from most other Tunisian ores, although it resem-bles the setting of Pb–Znmineralization that has been found inthe Oued El Abeid deposit in Central Tunisia (Fuchs 1973).

Following the discovery of Sekarna in 1906 (Sainfeld1952), the area was explored and then mined from 1922 to1948. Total production during this time was 4,540 tonnes ofgalena (2,600 t of Pb metal) and 22,000 tonnes of calamine(7,250 t of Zn metal). Mining activity focused on calamineand replacement galena ores in karstic cavities in Late Creta-ceous and Early Eocene carbonates. Renewed interest, whichhas focused on the Saint Eugène mine, stems from recentadvances in ore processing technologies and targets the non-sulfide Zn potential. Wide variation in the size of mineralizedzones at Sekarna complicates the estimation of tonnage andgrade for both the major and minor orebodies. Metal ratiosvary from non-sulfide Zn-rich ore, which has Zn/(Zn+Pb)>0.73, to sulfidic Pb–Zn ore, which has Zn/(Zn+Pb) of about0.35. Massive non-sulfide orebodies can reach 40,000 tonnes.

In this paper, we describe the tectonic and sedimentarysetting of the Sekarna mineralization, the mineralogy andtextures of the rocks, and the timing of mineralization rela-tive to sedimentation. We also report fluid inclusion datawhich constrain the temperature and salinity of the mineral-izing fluids, trace element data, and sulfur and lead isotopedata, which constrain the sources of sulfur and metals andthe mechanism of ore formation.

Geological setting

The Sekarna deposit is located in Central Tunisia, about220 km southwest of Tunis and 20 km SW of Maktar(Fig. 1). This area was located on the southern margin of theTethys Ocean during the LateMesozoic and Early Cenozoic, atime when most of Tunisia was submerged, except for Djef-fara Island to the south and Kasserine Island to the north (Sassi1974; Burollet and Odin 1980; Chaabani 1995; Zaïer et al.1998). The Sekarna area lay in the Northern Basin close toKasserine Island (Fig. 2), which has been subaerially exposedsince the Late Cretaceous (Sassi 1980; Béji-Sassi 1999; Zaïer1999). Kasserine Island controlled sedimentation in the sur-rounding basins, which included phosphatic sediments inshallow waters (Bonnefous and Bismuth 1982; Zaïer et al.1998). To the north, toward the continental margin, platformcarbonates (Nummulitic facies) and pelagic carbonates(Globogerines facies) were deposited.

The phosphorites of the Northern Basins are low in P205(<20%), fine-grained, compact, and glauconite-rich. They

normally comprise two phosphatic units separated by a thinmarl (Zaïer 1999), characteristically a nummulitic limestonebed.

Sekarna lies in a zone of Tertiary grabens (Graben Zone)that is bordered on the northwest by a zone of structuraldomes (Diapir Zone) and on the southeast by the CentralTunisian carbonate platform. Jebel (Mountain) Sekarnatrends NW–SE and lies east of the Rohia graben (Fig. 3a),midway between Jebel Serdj and Jebel Ajred. Jebel Sekarnawas affected by two systems of faults, the relative age ofwhich is difficult to determine with certainty. It is generallybelieved that NNW/SSE-oriented faulting preceded NNE–SSW faulting. The geometry of the Sekarna deposit wascontrolled by closely spaced syn-sedimentary fault systemsthat developed during NW–SE extension. These faults strikeNNW–SSE, NW–SE, or NNE–SSW and dip 40–60° to theW and NW. They appear mostly in the northern and westernparts of the study area and were responsible for lateral faciesvariations and thickness variations in the host sedimentaryrocks. The mineralized fault zones at Sekarna have approx-imately the same strike as the nearby Rohia graben (Fig. 3b).

The Sekarna area is underlain mainly by marine sedimen-tary rocks of Early Cretaceous to Eocene age (Fig. 4). Zaïer(1999) subdivided these rocks into five formations andsuggested that the Zn–Pb–(Ba) mineralization formed dur-ing diagenesis of some phosphorite zones (Fig. 5). The AlegFormation (Early Senonian) is a thick argillaceous marl withlimestone intercalations; it is a distinctive lithology thatoutcrops over wide areas and grades upward into the AbiodFormation (Late Senonian), an intensely fractured graymicritic limestone (30–50 m) with planktonic foraminferaand cherty lenses and nodules in its uppermost parts. Atother localities, the Abiod Formation contains upper andlower carbonate members; it is uncertain which member isrepresented at Sekarna.

The El Haria Formation (Paleocene) is too thin (2m thick) tobe shown on the map (Fig. 3b). It comprises black marls (illite–smectite mixed layers with minor kaolinite at the base and puresmectite at the top) with disseminated pyrite and glauconite,which implies marine regression and a suboxic–anoxic envi-ronment of deposition. This unit overlies the hard ground of theAbiod Formation and is overlain by transgressive conglomer-atic phosphorites of the lower Metlaoui Formation.

The phosphorite, which is contained within the MetlaouiFormation, is more lithologically diverse at Sekarna thanelsewhere in Tunisia (Gafsa Basin and Eastern and NorthernBasins) where it is hard, glauconite-rich, and well-silicified.The sequence consists of graded phosphatic–glauconiticmicroconglomerate beds (unit A) overlain by a thick well-bedded Early Eocene carbonate series with abundant num-mulites (units B and C).

Unit A is the richest in phosphorite and corresponds to the“Série Phosphatée.” It is a green to brown gray phosphatic–

546 Miner Deposita (2012) 47:545–562

glauconitic microconglomerate that outcrops as massivelayers (20–80 cm) with total thicknesses of 5 m (Saint Pierre)to 7 m (Ras El Guézir). The absence of open-marine fossilsand the abundance of clayeymud–wackstones suggest that theunit was deposited in restricted, poorly oxygenated deepwater, presumably during a transgression.

Unit B is the thickest member (20 m) of the MetlaouiFormation. Carbonate rocks with a conglomeratic appear-ance (“calcaires en boules”) alternate with dolomitic inter-beds. Unit C has at its base a 4- to 5-m-thick dark gray finedolomite with sporadic nummulites. These rocks are over-lain by 8–10 m of micritic limestone with nummulites. TheMetlaoui Formation is overlain by the Souar Formation, a

thick sequence dominated by marls of Late Eocene age.Sedimentary strata show facies changes and thicknesschanges over short distances. This is thought to reflectsyn-sedimentary faulting during Maastrichtian–Ypresiantectonic and/or halokinetic events.

Analytical methods

Representative samples of mineralized and non-mineralizedrocks were obtained from the Saint Pierre and Saint Eugènemines. All were examined by X-ray diffraction and trans-mitted and reflected light microscopy. Bulk rock chemical

Fig. 1 Map showing the location of the study area within the main tectono-sedimentary units of Tunisia and the locations of the major ore deposits(modified from Bouhlel 1993; Bouhlel et al. 2007; Perthuisot et al. 1987; Sainfeld 1952)

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analyses of six phosphorite samples were performed atActLabs (Ontario, Canada) using the 4LITHO, 4B-INNA,and 4B1 packages. Method descriptions can be found athttp://www.actlabs.com.

Trace elements in sphalerite and galena were analyzed atthe Department of Earth Sciences, Università della Calabria,Italy, by laser ablation–inductively coupled plasma massspectrometry (LA-ICP-MS) using an Elan DRCe (PerkinElmer/SCIEX) connected to a New Wave UP213 Nd-YAGlaser (213 nm). Ablation craters were 50 and 80 μm indiameter and were analyzed with a frequency of 10 Hzand effluence of about 20 J/cm2. Trace and rare earth ele-ment (REE) concentrations could be analyzed down to theparts per billion levels. External calibration was performedusing the NIST SRM 612 glass standard in conjunction withZn and Pb concentrations that had been determined inde-pendently by SEM-EDS. Calibrations were regularly

checked during analytical sessions using the BCR2 glassreference material. Raw data were processed using theGLITTER program (Barca et al. 2007).

Fluid inclusions in Saint Pierre sphalerite was examinedin 20 doubly polished sections. The microthermometry wasperformed on ∼200-μm-thick wafers using a LinkamTHMS-600 Heating Freezing stage. The stage was calibrat-ed between −180°C and 600°C by measuring phase changesin fluid inclusions of known composition. Measurementerror was ±2°C for homogenization temperatures and±0.2°C for melting temperatures. Salinities were determinedfrom the last melting temperature of ice using the equationof Bodnar (1993) for the H2O–NaCl system. Liquid/vaporphase ratios were estimated using standardized charts (e.g.,Shepherd et al. 1985) with an error of ±10 vol.%.

Sulfur isotope compositions were determined for 22 sulfideand sulfate samples at the USGeological Survey in Denver bythe continuous flow method (Giesemann et al. 1994) using aThermoFinnigan Delta mass spectrometer. Reproducibilitywas ±0.2‰. The data are reported as per mil deviations fromthe Vienna Cañon Diablo Troilite (VCDT) standard. Carbonand oxygen isotope compositions were determined by thephosphoric acid method (McCrea 1950). Mass spectrometrywas carried out using a Finnigan MAT 252. The results arereported as per mil deviations from the Vienna Pee DeeBelemnite (VPDB) and Vienna Standard Mean Ocean Water(VSMOW) standards, respectively. Reproducibility was betterthan ±0.2‰ (1σ) for both δ13C and δ18O.

For lead isotopic compositions, galena was separated bycareful handpicking under a binocular microscope and thenrinsed with doubly distilled water. Aliquots weighing about1 mg were analyzed by thermal ionization mass spectrom-etry at the Geochronology and Isotope Geochemistry Labo-ratory, University of North Carolina, Chapel Hill, followingthe procedure described by Skaggs (2010).

Host rock petrography and geochemistry

Petrography

The ore-hosting phosphorite is similar to other TunisianLate Paleocene–Early Eocene phosphorites (Sassi 1974;Chaabani 1995; Béji-Sassi 1999; Zaïer 1999). The rockfabrics and lithostratigraphic position are consistent with acommonly invoked model that involves authigenic mineralgrowth in fine-grained marine sediments located at the oxic–suboxic interface in nearshore depositional environments(Baturin 1982; Brunett et al. 1983; Froelich et al. 1988;Föllmi 1990; Jarvis 1992; Krajewski et al. 1994).

The phosphorites are granular and include authigenicphosphatic grains (pellets, coprolites, phosphatized fossils);biogenic grains (fossil teeth, skeletal fragments); non-

Fig. 2 Palaeogeographic setting of Tunisia during the Early Eoceneshowing the location of the study area (modified from Sassi 1974;Winnock 1980; Zaïer et al. 1998)

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phosphatic grains (quartz, glauconite, lithoclast, microcrys-talline carbonate aggregates); and cement (carbonates, sili-cates) with packstone to grainstone textures. The phosphaticgrains are dominated by carbonate fluorapatite (CFA). Theclay fraction of the phosphorite, which reaches 1–5%, iscomposed of kaolinite, glauconite, and minor illite.

Carbonates dominate the cement in non-mineralizedphosphorite. In contrast, mineralized rocks are depleted incarbonate and enriched in silica. Silicified rocks are charac-terized by euhedral quartz, particularly in mineralized areas,which occurs as replacements of carbonate cement and asinclusions within sphalerite and barite. Quartz in some sam-ples shows growth layers and birefringent inclusions of

anhydrite. Unlike phosphorite and glauconite grains, whichshow sutured flat grain shapes, concavo-convex contacts,and other textural signs of compaction, the euhedral quartzgrains are undeformed, displaying euhedral shapes and lack-ing strained extinction. This implies that silicification post-dated compaction. Silicification was commonly intense,filling open spaces and locally replacing the carbonate ce-ment with euhedral quartz crystals with straight intercrys-talline boundaries.

Silicification appears from textural observations to haveoccurred relatively late in the diagenetic history of thephosphorites. It is linked to the base metal sulfide mineral-ization, but likely represents a pre-mineralization stage.

Fig. 3 a Regional geological map showing the Rohia graben (modified from Jauzein 1967). b Geological map of the study area showing thelocation of the main deposits (modified from Zaïer 1999)

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Geochemistry

Whole-rock chemical analyses (Table 1) show that mineral-ized rocks are dominated by SiO2 (30.3–39.8%), CaO, andP2O5, with all other oxides (MnO, MgO, Na2O, K2O, TiO2)below about 1%. Non-mineralized samples exhibit higherCaO (38.1–42.3%) and MgO (2.4–5.1%). P2O5 ranges from19.1% to 23.7% and is slightly higher in non-mineralizedsamples. Al2O3, TiO2, and K2O are indistinguishable be-tween mineralized and non-mineralized rocks, reflectingrelatively uniform contributions of detrital material.

Zn, Pb, and Ba in mineralized samples (n02) average3,315, 1,520, and 4,269 ppm, respectively. The highestgrade samples contained 5.3% Zn and 2.6% Pb. Minor totrace As, Ag, and Sb substitute for Pb in the galena lattice,and Cd substitutes for Zn, giving rise to relatively highconcentrations of these elements in mineralized samples.Similarly, Sr substitution in the barite lattice explains theelevated Sr in mineralized samples. Non-mineralized phos-phorites have higher Ni and Mo (70–90 and 11–31 ppm,respectively) than mineralized samples (20–30 and 11–16 ppm, respectively). Ni and Mo are positively correlatedwith P2O5 and Na2O, which indicates that these metals arerelated to CFA minerals.

Compared with phosphorite from the Gafsa–MetlaouiBasin (sample KEC1), non-mineralized phosphorite atSekarna is slightly enriched in V, Ba, Ni, Mo, U, and∑REE and is depleted in Sr, Zn, and Cd. Variations in thetrace elements of phosphorite in the two localities can beattributed either to varying depositional conditions (basinconfiguration, redox conditions, productivity) or mixing ofchemical precipitates and detritus shed from chemicallydistinct source areas.

Fig. 4 Outcrops of Late Cretaceous–Early Eocene sedimentarysequences in Ras El Guezir (1,322 m), 800 m from the Saint Pierredeposit. The microconglomeratic phosphorite lies between reducedmarls (∼1.5 m thick) of El Haria Formation and the upper carbonate(∼30 m thick) of Metlaoui Formation

Fig. 5 Simplified composite lithostratigraphic column for the Sekarnaarea showing the stratigraphic position of the major sulfide and non-sulfide deposits

b

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∑REE and Y contents of the samples average 741.5 and295 ppm, respectively, higher than average phosphorite(462 ppm REE, 275 ppm Y; Altschuler 1980). PAAS-normalized (McLennan 1989) REE patterns display a slightenrichment in heavy REE (HREE), with (La/Yb)N of 0.85–0.94 and (La/Sm)N of 0.69–0.86. All samples have high (La/Yb)N ratios relative to modern seawater (0.2–0.5; Reynardet al. 1999), probably reflecting adsorption onto the apatitelattice during early diagenesis. The yttrium anomaly in oursamples varies from 1.54 to 1.66, and the La/Nd ratio variesfrom 0.99 to 1.19, which is indistinguishable from seawater(Shields and Stille 2001).

Sulfide minerals are very low in ∑REE (<3 ppm) com-pared with CFA, so their contribution to whole-rock REEcontent is negligible. The good correlation between P2O5

and ∑REE (R200.72) suggests that the REE inventory ofthe rocks is controlled by the phosphate components. Non-mineralized and mineralized phosphorites display similarREE patterns (Fig. 6); both resemble seawater, with nega-tive Ce anomalies and flat or slight enrichment in HREEenrichment. It would appear that both phosphorite typesobtained REE from seawater and record depositional con-ditions similar to modern oxic–suboxic seawater.

The Sekarna patterns resemble patterns that have beenreported for phosphorites elsewhere in Tunisia (Béji-Sassi1999; Ounis et al. 2008) as well as phosphorites of Meso-zoic and Cenozoic age elsewhere in the Tethyan province(McArthur and Walsh 1984).

Mineralization

Saint Pierre deposit

Both hypogene and supergene Zn–Pb deposits occur inSekarna area. The Saint Pierre deposit is a hypogene strat-iform Zn–Pb–(Ba) sulfide deposit (10–13% Zn, 3–5% Pb)that is hosted by unusually hard, massive, dark gray, phos-phatic–glauconitic microconglomerate of Ypresian age(Fig. 7a). Additional massive galena (3–14%) and super-gene non-sulfide Zn minerals, such as Fe-poor smithsonite(<2 wt.% FeO), hemimorphite, and hydrozincite, occur inthe Late Cretaceous carbonate rocks in karstic cavities alongN20°–N40° trending faults. The mineralized unit is un-derlain by Paleocene marls and overlain by Nummliticlimestones. The mineralization appears to fill a strati-form lens-shaped structure concordant to bedding andtrending WNW to NW. The mineralized zones areassociated with major faults that are interpreted to havebeen feeder channels for hydrothermal fluids. The min-eralization extends 200–300 m along strike and 100 macross strike. The sulfides do not appear to haveformed by filling open space in fractures. It is likely

that the mineralization was emplaced after the mostsignificant folding and faulting events that affected thearea. Mineralization is texturally and mineralogicallysimple, comprising sphalerite (90% of the ore), galena,Fe sulfides, and barite, and was preceded by a silicification ofthe host rock carbonate. Sphalerite and galena occur ascements around phosphatic pellets, glauconite, and detritalor euhedral quartz (Fig. 8a–d).

The ore-stage sulfides formed after burial diagenesis bythe dissolution and replacement of either the rock matrix orearlier carbonate cements. Sulfide textures clearly indicatethat the minerals grew during the late diagenetic to epige-netic stages in the evolution of the host sequence. Barite,which averages ∼5.76% Sr, is a minor constituent of silici-fied phosphatic marls where it occurs as fibroradiated con-cretions. The concretions are up to 3 cm in diameter,spherical to ellipsoidal in shape, radial in structure, and arecomposed of elongate barite prisms and fibers. Barite isintergrown with phosphorite grains, euhedral quartz, andglauconite. Barite also fills intergranular space and micro-fractures in some phosphatic grains and occurs as microm-eter inclusions in sphalerite (Fig. 8e, f).

Sphalerite, the major ore sulfide, occurs as disseminatedgrains or massive aggregates that are typically reddishbrown to yellow, anhedral to subhedral, and medium-grained (1–10 mm). In transmitted light, some sphaleritecrystals are finely zoned, showing alternating reddish andyellow zones and distinct growth layers. Galena occurs aslarge (>2 cm), well-formed crystals in simple cubic formsand, like sphalerite, as dendritic intergrowths. Galenaappears to have formed late in the paragenetic sequence asit commonly cements other sulfides. Massive galena occurswith calamine in the Abiod and Upper Metlaoui carbonaterocks. Colloform cerussite and anglesite are intergrownwithin fractures in galena and along grain boundaries.

Iron sulfides occur as minor disseminations in all parage-netic stages. Pyrite formed early in the paragenesis (diageneticpyrite), whereas later paragenetic stages are characterized bymixed marcasite and pyrite. The earliest pyrite typicallyoccurs as framboids or micron-sized euhedral crystals dissem-inated in the argillaceous matrix or between larger mineralgrains. Pyrite also occurs as rims and as replacements ofglauconitic and phosphatic pellets, foraminifers, and gastro-pods. The upper part of the Saint Pierre ore body is oxidizedand contains supergene minerals, including Cd-rich smithson-ite, Zn silicates, Fe oxyhydroxides, and minor Pb minerals.

Saint Eugène deposit

Saint Eugène is a non-sulfide Zn–Pb deposit (35–40% Zn).The ore bodies are hosted by dolomitic limestone and areassociated with two major NE/SW-trending faults. The mainZn mineral is smithsonite. Hemimorphite, hydrozincite,

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cerussite, and anglesite are minor constituents of the deposit(Fig. 7b). Calcite is often associated with smithsonite. Feoxyhydroxides are also very common.

Smithsonite δ18O and δ13C values are −5.1‰ and28.3‰, respectively. Dolomitic host rocks show a relativelynarrow range of δ18O (21.6–23.9‰), and a wider range ofδ13C (−2.3‰ to 2.6‰, Table 2). The dolomite δ13C valuessuggest that carbon was derived from both dissolution of theprecursor limestones and oxidation of organic carbon.

The oxygen isotopic compositions of Sekarna smithson-ites are similar to the compositions that have been reportedfor other supergene smithsonites elsewhere in the world(Boni et al. 2003; Coppola et al. 2008, 2009; Gilg et al.2008). Both the mineral assemblage (Fe-poor smithsonite,Fe oxyhydroxides, Pb sulfate carbonates) and the C–Oisotopic signature are typical of ores of supergene origin(Gilg and Boni 2004a, b; Gilg et al. 2008). The fact thatdolomite δ18O values are about 4‰ lower than the smith-sonite values is in accord with the model of isotope frac-tionation between carbonate minerals and water proposedby Gilg et al. (2008). The fluid that oxidized the precursorsulfides at Saint Eugène was likely of meteoric origin.Carbonate block faulting during the late stages of the EarlyTertiary Alpine orogeny could have enhanced deep circula-tion of meteoric fluids and the oxidation of primary sulfides.

Table 1 Major (%) and trace element (in parts per million) contents ofrepresentative phosphorite samples

Sample SEK1 SEK3 SEK4 SEK5 SEK2 KEC1NMP NMP NMP MP MP GM

SiO2 (%) 18.20 22.45 9.31 30.38 39.88 6.97

Al2O3 1.22 1.21 0.88 1.51 1.47 1.28

Fe2O3(T) 2.17 1.99 1.69 1.75 1.68 0.56

MnO 0.03 0.02 0.02 0.002 0.003 0.003

MgO 3.88 2.43 5.12 0.37 0.42 0.62

CaO 38.76 38.18 42.35 32.61 28.05 45.88

Na2O 0.41 0.48 0.42 0.42 0.38 1.42

K2O 0.50 0.45 0.39 0.61 0.54 0.32

TiO2 0.04 0.05 0.04 0.07 0.07 0.053

P2O5 19.73 23.77 19.14 23.66 20.15 27.81

LOI 13.74 9.06 19.02 5.60 4.90 11.9

Total 98.67 100.10 98.38 96.99 97.56 96.81

V (ppm) 254 164 183 297 275 52

Ba 45 99 89 4,563 3,975 40

Sr 539 573 693 851 785 1,816

Y 299 316 251 326 284 140

Zr 34 41 28 41 41 44

Cr 297 305 287 482 437 273

Ni 64 85 58 31 21 22

Cu 15 16 13 21 34 11

Zn 84 33 13 3,400 3,230 290

Cd 5.6 2.9 0.8 349 302 60.8

As 18.7 37.2 37.6 68.2 107 6.5

Rb 12 11 9 15 13 6

Nb 1 2 1 3 3 2

Mo 11 31 25 16 16 11

Ag <0.3 1.1 0.8 2 2 0.9

Sb 2.8 4.9 5.4 8 45.3 1

Pb 31 376 207 1,270 1,770 5

Th 9.3 9.7 5.5 7.5 6.8 14.1

U 36.2 55 47.8 67.3 57.2 34.2

La 184 186 155 197 166 108

Ce 195 213 152 223 182 164

Pr 38.90 41.90 30.50 45 38.20 23.5

Nd 167 182 130 197 167 96.5

Sm 34.70 37.60 26.20 41.40 35 18.9

Eu 8.26 9 6.41 9.93 8.32 4.34

Gd 37 40.10 29 43.50 36.50 18.4

Tb 5.60 6 4.40 6.40 5.50 2.8

Dy 32.50 34.50 25.50 36.60 31 16.1

Ho 7.10 7.50 5.70 7.80 6.60 3.5

Er 19.60 20.80 15.90 20.80 17.90 10.1

Tm 2.56 2.67 2.09 2.64 2.26 1.38

Yb 15.80 16.10 12.80 15.40 13.10 9.2

Lu 2.39 2.46 2 2.28 1.98 1.46

∑ REE 750.41 799.63 597.50 848.75 711.36 478.18

Ce/Ce* 0.50 0.53 0.48 0.52 0.50 0.73

Eu/Eu* 2.03 2.12 1.79 2.24 2.05 1.48

(La/Sm)N 0.77 0.72 0.86 0.69 0.69 0.83

Table 1 (continued)

Sample SEK1 SEK3 SEK4 SEK5 SEK2 KEC1NMP NMP NMP MP MP GM

(La/Yb)N 0.86 0.85 0.89 0.94 0.94 0.87

(Dy/Yb)N 1.24 1.29 1.20 1.43 1.43 1.05

Y ano 1.57 1.57 1.66 1.54 1.58 1.49

La/Nd 1.10 1.02 1.19 1.00 0.99 1.12

LOI loss on ignition, NMP non-mineralized phosphorite, MP mineral-ized phosphorite, GM Gafsa–Metlaoui phosphorite

Ce/Ce*=3CeN/(2LaN+NdN) (The subscript N refers to normalizationof concentrations against the PAAS values)

Eu/Eu*=EuN/(SmN+GdN)0.5

0,1

1

10

100

La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

RE

E /P

AA

S

SEK1

SEK3

SEK4

SEK5

SEK2

KEC1

Fig. 6 PAAS-normalized REE patterns of mineralized and non-mineralized phosphorites. PAAS composition after McLennan (1989)

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Fig. 7 a Typical ore sample from Saint Pierre. Fine-grained sphalerite(reddish to yellow) has replaced carbonate cement in Eocene phosphorite(dark color). b Typical ore sample from Saint Eugène concretionary,colloform calamine ore in karstic cavities within Early Eocene limestone.

1 Greenish to yellow colloform and concretionary smithsonite. 2 Whitecolloform hydrozincite. 3 Fe oxides with Fe-rich smithsonite. 4 Weath-ered carbonate host rock

Fig. 8 a–d Photomicrographsof polished sections illustratingsulfide–host rock textures. aPolarized reflected lightphotomicrograph showingyellow red sphalerite (sp)cementing phosphatic pellets(ph). b Polarized reflected lightphotomicrograph showingsphalerite infilling betweenphosphatic pellet and detritalquartz (qz). Note the tangentialand sutured contacts betweensphalerite, detrital quartz, andphosphorite grains, indicativeof intergranular pressure. cSphalerite cementing euhedralquartz and phosphorite grains. dSphalerite surrounding detritalquartz. e, f Backscatteredelectron images showing therelationship between barite andthe surrounding minerals. eBarite inclusions (white spots)in sphalerite in a matrix offrancolite. Francolite containsup to 1.8 at.% Na and up to4.6 at.% F. Cd in sphalerite isup to 0.8 at.% (equivalent to 1.9w.t% Cd). f Fibrous baritereplacing carbonate cement andshowing irregular outlines withrounded to sub-rounded grains

Miner Deposita (2012) 47:545–562 553

During this period, several short-lasting emersions and in-tense weathering episodes facilitated karst formation in theLate Cretaceous and Early Eocene carbonate rocks andoxidation of sulfide bodies by downward-percolating mete-oric waters.

Geochemistry

Trace elements

Data from nine polished sections (Tables 3 and 4) indicatethat Sekarna sphalerite is low in Fe (average 652 ppm) andhigh in Cd (average 11,337 ppm). Cadmium concentrationsare variable, ranging from 5,789 to 20,106 ppm (Fig. 9).Within an individual hand specimen, Cd varies by as muchas a factor of 3. Fe and Cu contents range from 344 to1,312 ppm and from 25 to 1,207 ppm, respectively. Withinhad specimens, Fe and Cu vary by a factor of <3.

Sekarna sphalerite comprises a single generation thatdisplays color-banded growth zones corresponding to oscil-lations in minor element concentrations. LA-ICP-MS steptraverses show that light yellow zones in crystal cores arelower in Cu and higher in Cd and Fe than reddish brownzones in crystal rims. The increase in Cu toward crystal rimsmay reflect the introduction of Cu-rich fluids.

Neither Cu- nor Cd-bearing minerals were present asinclusions in sphalerite; thus, these two elements wouldappear to be in solid solution in the sphalerite lattice. Incontrast, submicroscopic pyrite was identified in some spha-lerites. Strong correlations in spot analyses between As andPb, As and Sb, and Pb and Sb (R200.6) could reflectsubmicroscopic inclusions of galena and/or tetrahedrite.

Trace elements in Sekarna galena are predominantly As, Ag,Cd, and Sb. Sb ranges up to 1,740 ppm and Ag from 4 to122 ppm. Cd is higher than in most Mississippi Valley-type(MVT) deposits (Hall andHeyl 1968; Hagni 1983; Song 1984).

Elevated Fe, Cu, and Zn in galena were almost certainlydue to the ablation of microscopic inclusions of pyrite, chal-copyrite, and sphalerite. Strong correlations between As andCd (R2 above 0.8) suggest that these elements are probablyassociated with inclusions of minerals such as sphalerite.

Fluid inclusions

Fluid inclusions in Saint Pierre sphalerite are 20–80 μm indiameter. According to their appearance at room tempera-ture, two types of inclusions were recognized: single-phaseliquid aqueous inclusions (L) that are abundant and gener-ally smaller than 20 μm and two-phase inclusions (L+V)that are located along the growth zones and are about 90%liquid. Inclusion appearance is highly variable and can becharacterized as irregular, slightly rounded and flattened,tubular, or spheroidal/oblate.

Homogenization temperatures (TH) were determined princi-pally on two-phase liquid-rich inclusions, which were relativelylarge and regular in shape. These inclusions showed no texturalevidence of necking down and are interpreted to be primary.

Microthermometric results ranged from 80°C to 130°C.First ice melting temperatures ranged from −59°C to −34°C.These temperatures, which are lower than the eutectic tem-perature in the NaCl–H2O system (Te0−21.2°C), indicatethat the fluid is complex and may contain Ca2+ (H2O–NaCl–CaCl2: Te≈−55°C) and/or Mg2+ (H2O-NaCl-MgCl2: Te≈−35°C; Crawford 1981; Roedder 1984; Shepherd et al.1985). Final ice melting temperatures ranged from −11°Cto −22°C, which indicates a salinity range of 15–24 wt.%NaCl eq. The measured salinities are lower than the eutecticsalinity for the binary system NaCl–H2O (23.3 wt.%NaCl eq).

The fluid inclusion observations allowed two types offluids to be identified in Sekarna sphalerite:

1. An H2O–NaCl–CaCl2, moderate temperature (110–130°C) and moderate salinity (∼15 wt.% NaCl eq) fluidthat is predominant in brown reddish zones of sphalerite

2. An H2O–NaCl–CaCl2, moderate- to low-temperature(80–110°C) and high-salinity (∼24 wt.% NaCl eq) fluidthat is associated mainly with light yellow crystal cores

The measured temperatures and calculated salinities liewithin the ranges that have been reported for MVT depositsworldwide (50–250°C, 10–30 wt.% NaCl eq). The variationsin salinity and eutectic temperature in primary two-phase fluidinclusions in zoned sphalerite could reflect mixing betweentwo chemically distinct fluids or phase separation of a singlefluid.

Table 2 Carbon and oxygen isotopic compositions of smithsonite and dolomite

Sample references Composition δ13CVPDB (‰) δ18OVSMOW (‰)

SM1 Yellow colloform smithsonite from Saint Eugène −5.1 28.3

SM2 Massive crystalline smithsonite from Saint Eugène −5.2 28.3

D1 Early Eocene massive gray dolomite from Saint Eugène (unit C) hosted sample SM1 −2.3 23.9

D2 Early Eocene yellow dolomite from Saint Eugène (Unit C) hosted sample SM2 2.3 21.6

A1 Abiod dolomitized limestone from Saint Pierre 2.6 23.0

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Sulfur isotopes

The sulfur isotopic compositions (δ34SVCDT) of ore sulfidesat Sekarna are relatively uniform with averages of −10.2‰for sphalerite and −13.6‰ for galena (Table 5). Disseminat-ed sphalerite is isotopically indistinguishable from massivesphalerite. Disseminated galena within phosphorite(−12.5‰, n05) is slightly heavier than massive galena(−15.5‰, n03) in the Late Metlaoui carbonate rocks(Fig. 10). Most of the δ34S results for sulfate minerals(barite, celestite, and gypsum) range from 17.7‰ to24.9‰, values that are similar to marine sulfates (e.g.,Claypool et al. 1980).

Lead isotope compositions

The Pb isotope results for seven Sekarna galenas are com-pared in Fig. 11 with the Pb isotopic compositions of themajor Pb–Zn Tunisian deposits (Bougrine, Fedj El Adoum,and Boujabeur) as well as the Pb isotope evolution curves ofZartman and Doe (1981). 206Pb/204Pb ranges from 18.7418to 18.8627, 207Pb/204Pb from 15.6488 to 15.7902, and208Pb/204Pb from 38.7976 to 39.27. Sekarna galena is

isotopically homogenous, plotting in narrow fields that im-ply an upper crustal Pb source.

The Pb isotopic compositions of the major Tunisian oredeposits do not show a trend representing enrichment inradiogenic lead. There is no systematic relationship betweenPb isotopes and geographic location, nor is there a relation-ship between Pb isotopes and age of wall rocks or deposittype. All the galenas plot along the upper crustal trend. Theisotopic uniformity indicates a Pb source that was well-homogenized (Table 6).

Discussion

Diagenetic vs. epigenetic origin

At Saint Pierre, the Zn–Pb–(Ba) mineralized zone occurswithin the lower Metlaoui organic-rich unit. The depositwas emplaced in an extensional zone at the edge of theRohia Graben, the boundary fault of which has the samestrike as the mineralized fault zones at Sekarna. The miner-alized zone corresponds to a local depocenter immediatelyadjacent to faults that provided conduits for the ore fluids.

Table 3 LA-ICP-MS analyses of sphalerite from Saint Pierre

Sample references Spot analyses Description Ti Mn Fe Co Cu As Ag Cd Sn Sb Pb

P-SP (disseminated sphalerite) 03-P-SP-01 Brown reddish sphalerite 5 nd 451 0.5 174 nd 572 5,789 0.1 1 28

04-P-SP-02 Brown reddish sphalerite 4 nd 344 0.5 1,019 6 31 6,822 nd 3 20

05-P-SP-03 Brown- reddish sphalerite 6 0.4 366 0.5 836 4 77 7,335 0.2 2 20

06-P-SP-04 Light yellow sphalerite 7 0.2 480 0.3 593 46 4 11,705 0.1 51 63

07-P-SP-05 Brown reddish sphalerite 3 nd 388 0.4 1207 72 6 6,528 nd 118 92

08-P-SP-06 Light yellow sphalerite 3 0.2 461 0.3 227 41 4 14,230 0.1 26 33

09-P-SP-07 Light yellow sphalerite 7 nd 531 0.3 51 2 nd 11,882 0.2 4 7

10-P-SP-08 Light yellow sphalerite nd nd 419 0.5 215 21 3 12,654 nd 22 24

SP1 (disseminated sphalerite) 08-SP1-01 Light yellow sphalerite 2 0.2 676 0.4 765 70 10 16,001 1.4 144 129

09-SP1-02 Brown reddish sphalerite 1 nd 662 0.7 368 41 8 10,347 2.0 72 40

11-SP1-04 Light yellow sphalerite nd 0.4 967 0.6 529 54 13 16,272 4.1 145 74

SP3 (massive sphalerite) 09-SP3-01 Light yellow sphalerite 2 nd 852 0.4 188 32 11 14,845 nd 68 46

10-SP3-02 Brown reddish sphalerite 1 0.2 686 0.3 131 28 10 9,578 0.1 74 42

11-SP3-03 Light yellow sphalerite 2 nd 571 0.3 398 21 6 12,399 nd 37 87

13-SP3-05 Brown reddish sphalerite 2 0.3 628 0.3 415 49 13 11,844 0.1 75 109

SP4 (massive sphalerite) 21-SP4-01 Brown reddish sphalerite nd nd 600 0.5 774 57 7 9,158 2.7 38 69

22-SP4-02 Light yellow sphalerite 2 0.1 596 0.4 820 50 7 10,120 0.1 94 97

SP-G4 (massive sphalerite) 19-SP-G4-01 Light yellow sphalerite 3 0.3 1,312 0.3 360 13 4 16,158 0.1 40 50

20-SP-G4-02 Brown reddish sphalerite 2 0.1 914 0.4 422 11 4 10,243 0.1 61 37

SP-G5 (massive sphalerite) 15-SP-G5-01 Light yellow sphalerite 1 0.1 622 0.2 25 3 1 12,028 nd 9 34

16-SP-G5-02 Brown reddish sphalerite 1 0.1 579 0.2 138 22 9 7,721 0.1 51 24

17-SP-G5-04 Light yellow sphalerite 1 0.2 1,183 0.4 119 7 3 20,106 nd 26 16

18-SP-G5-03 Brown reddish sphalerite 1 0.1 712 0.3 1,072 25 9 6,988 0.5 148 86

Trace element concentrations in parts per million

nd not detected

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The mineralization does not fill fractures but rather islens-shaped and interstratified with phosphorite. Thefinely disseminated nature of the sulfides suggests thatsulfide minerals were deposited where favorable horizons inthe phosphorite were dissolved. The spatial correspondence ofsulfide lenses and faults, the depositional temperatures rang-ing from 80°C to 130°C, and the homogeneous S and Pb

isotopic compositions support a model in which fluidswere expelled from a proximal depocenter by sedimentcompaction.

The morphology, stratigraphic position, and geochemis-try of REE, in particular of the phosphorites host rocks, aresimilar to the unmineralized CFA in other localities thatformed during the diagenesis of organic-rich sediments.For this reason, the phosphorites are regarded as primary.The absence of diagenetic intergrowths of sphalerite orgalena with phosphorite and glauconite suggests that sulfidedeposition postdated diagenesis. Although carbonate ce-ment is preserved in non-mineralized phosphorites, silicifi-cation and sulfide growth have destroyed sedimentary andearlier diagenetic textures within the ore zones. Sulfides didnot replace phosphatic minerals; instead, sphalerite and ga-lena formed as cements around phosphatic pellets, glauco-nite, and quartz (both detrital and euhedral). Taken together,these observations suggest a late diagenetic to epigeneticorigin for the Zn–Pb mineralization.

Sphalerite in the Sekarna deposit is characteristically fineand reddish yellow in color with significant enrichment inCd (average011,337 ppm), some 21 to 74 times richer thansphalerite in other major MVT deposits in the world. HighCd, similar to Sekarna, is reported in a few MVT deposits,

Table 4 LA-ICP-MS analyses of galena from Saint Pierre and Saint Eugène

Sample references Spot analyses Ti Fe Cu Ni Zn As Ag Cd Sb

P-G (disseminated galena in phosphorites from Saint Pierre) 03-P-G-01 0.5 nd nd nd 0.1 3.3 44 66 296

04-P-G-02 0.7 nd nd 0.1 0.7 0.9 105 28 239

05-P-G-03 1.1 7.8 0.1 0.3 0.2 1.0 82 27 581

06-P-G-04 0.7 nd nd 0.1 nd 2.7 106 37 295

07-P-G-05 1.5 5.1 0.5 0.2 4.5 0.4 105 29 451

08-P-G-06 2.7 9.7 0.1 0.2 nd nd 37 12 68

SP-G1 (massive galena in phosphorites from Saint Pierre) 03-SP-G1-01 nd nd 1.1 nd 0.2 4.0 29 77 846

04-SP-G1-02 0.9 nd 1.7 nd nd 2.6 38 50 679

05-SP-G1-03 0.2 nd 0.2 0.1 0.3 0.4 9 25 610

SP-G2 (massive galena in phosphorites from Saint Pierre) 06-SP-G2-01 0.7 nd nd nd nd 0.5 5 21 239

07-SP-G2-02 0.6 nd 0.1 nd nd 5.7 8 88 233

08-SP-G2-03 0.5 nd nd nd 0.2 5.5 6 82 237

09-SP-G2-04 0.4 nd nd 0.0 nd 1.9 8 42 219

10-SP-G2-05 0.8 nd nd nd nd 0.2 4 24 68

SP-G3 (massive galena from Saint Eugène) 11-SP-G3-01 1.0 nd 1.8 0.1 0.4 1.3 44 38 182

12-SP-G3-02 0.8 nd 0.1 nd nd nd 38 23 84

13-SP-G3-03 0.8 1.5 0.1 nd 0.2 nd 26 19 70

14-SP-G3-04 0.2 0.9 1.1 0.1 0.1 6.8 64 108 263

G2 (massive galena from Saint Eugène) 10-G2-01 1.1 nd 0.8 nd 0.2 1.5 15 37 1,468

12-G2-03 1.1 3.8 1.5 0.1 0.5 0.4 122 59 1,309

14-G2-05 1.1 nd 0.1 nd 0.7 1.1 27 92 1,740

G3 (massive galena from Saint Eugène) 06-G3-01 0.3 nd 0.3 0.1 nd 3.1 28 92 363

09-G3-04 1.3 nd 0.3 nd nd 15.3 26 189 675

Trace element concentrations in ppm

nd not detected

1

10

100

1000

10000

1 10 100 1000 10000 100000

Sb (p

pm)

Cd (ppm)

Sphalerite

Galena

Fig. 9 Cd vs. Sb plot of laser ablation ICP-MS data for sphalerite andgalena from Sekarna

556 Miner Deposita (2012) 47:545–562

including Kentucky–Tennessee (average013,344 ppm; Jollyand Heyl 1968), Vazante (average08,410 ppm; Monteiro etal. 2006), and the low-temperature Cd-rich zinc deposit of

Niujiaotang, China (average013,800 ppm; Liu et al. 1999);high Cd sphalerite is also known in phosphate-rich sediments(Nathan et al. 1996, 1997; Béji-Sassi and Sassi 1999).

Table 5 Sulfur isotopic compositions of sphalerite, galena, and barite from the Sekarna deposit and of celestite and gypsum from southern Tunisia

Samplereferences

Location Composition Mineralization type δ34S (‰)

SEKZn1 Saint Pierre Sphalerite Disseminated in phosphorites −9.5

SEKZn2 Saint Pierre Sphalerite Disseminated in phosphorites −9.3

P-SP/SP Saint Pierre Sphalerite Disseminated in phosphorites −10.8

SP1 Saint Pierre Sphalerite Disseminated in phosphorites −10.1

SP2 Saint Pierre Sphalerite Disseminated in phosphorites −10.3

SP3 Saint Pierre Sphalerite Disseminated in phosphorites −10

SP4 Saint Pierre Sphalerite Disseminated in phosphorites −10.8

SP5 Saint Pierre Sphalerite Massive ore −11.2

SP6 Saint Pierre Sphalerite Massive ore −10

SEKPb1 Saint Pierre Galena Disseminated in phosphorites −12.8

SEKPb2 Saint Pierre Galena Disseminated in phosphorites −12.9

GN1 Saint Pierre Galena Disseminated in phosphorites −12.3

GN2 Saint Eugène Galena Massive ore in dolomite −16

GN3 Saint Eugène Galena Massive ore in dolomite −15.5

GN4 Saint Eugène Galena Massive ore in dolomite −15.1

PG-G Saint Pierre Galena Disseminated in phosphorites −12.4

Ba1 Saint Pierre Barite Fibroradiated barite in phosphorites 25.3

Ba2 Saint Pierre Barite Fibroradiated barite in phosphorites 24.9

CMK1 Jebel Jebes (Eastern basins) Nodular celestite Celestite-bearing dolostone (El Haria Formation) 21.2

C-GAF Jebel Chouabine (Gafsa–Metlaouibasin)

Celestite geode Celestite-bearing evaporitic unit of Thelja member (Paleocene–EarlyEocene)

22.5

G-CELS Jebel Jebes (Eastern basins) Gypsum Early Eocene gypsum associated to nodular celestite (sample CMK1) 17.8

G-TH Oued Thelja (Gafsa–Metlaoui Basin) Gypsum Eocene gypsum in massive gypsum deposit (Thelja member) 17.7

G-JJB Jebel Jebes (Eastern basins) Gypsum Gypsum filling fractures in phosphorite beds −9.5

δ34S (VCDT)

-15 -10 -5 0 5 10 15 20 25

Fre

quen

cy

0

1

2

3

4

5

Disseminated sphalerite (Saint Pierre)Massive sphalerite (Saint Pierre)Disseminated galena (Saint Pierre)Massive galena (Saint Eugène)Barite (Saint Pierre)Celestite (Southern Tunisia)Eocene gypsum (Southern Tunisia)

Fig. 10 Frequency distributionof S isotope values (per mil) forsphalerite, galena, and barite inthe Sekarna deposit

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The high Cd suggests that the hydrothermal fluid mayhave circulated through the underlying basin fill, whichcontains organic-rich sediments. Similarities in sphaleritecomposition between the Sekarna deposits and low-temperature deposits elsewhere, including the MVT deposits,may be explained by brines from similar sources traversingand leaching metals from sedimentary sequences.

The TH and salinities of Sekarna sphalerite-hosted inclu-sions are consistent with an ore genesis model involvingbasinal brines like those that formed Tunisian MVT deposits(Bouhlel 2005) and MVT deposits in other areas of theworld (Leach et al. 2005). The relatively uniform Pb isoto-pic compositions of Sekarna galena suggests a simple originfor the metals, namely, derivation from a single fluid gen-eration that interacted with deep upper crust.

Diagenetic barite

Although sulfates (celestite, gypsum) have been describedpreviously in phosphorites of Tunisia (Béji-Sassi 1999),Sekarna is the first reported occurrence of concretionarybarite. Other stratigraphic intervals are known to containabundant barite concretions, including mid-Cretaceous sedi-ments, especially those of Albian age (France: Bréhéret andBrumsack 2000; Belgium: Dejonghe et al. 1987; Tunisia:Burollet et al. 1983), at the Cretaceous–Tertiary boundary(Soliman 1998; Ramkumar et al. 2005) and across thePaleocene–Eocene Thermal Maximum (PETM) period(Schmitz et al. 1997; Bains et al. 2000). Ba anomalies atthe Cretaceous–Tertiary have been attributed to a detritalinflux coincident with a fall in sea level (Ramkumar et al.2005). Similar anomalies across the PETM were attributedto increased productivity (Bains et al. 2000).

The close association of celestite with evaporates is at-tributed to sulfate enrichment in evaporated pore waters orby the dissolution of evaporites, or else as a consequence ofthe transformation of aragonite to calcite and/or dolomiteduring early diagenesis (e.g., Yan and Carlson 2003). Noevaporite sequences have been observed in Sekarna area,nor have evaporates been recognized in the sedimentaryrecord at Sekarna. Thus, evaporation processes are unlikelyto have been for the ultimate cause of celestite precipitation.

The sulfur isotopic compositions of celestite are consis-tent with literature data in showing a 5–10‰ enrichment inδ34S relative to associated gypsum (Table 5). The largeisotopic difference would seem to preclude simple dissolu-tion of gypsum and reprecipitation of the sulfate as celestite.Two mechanisms could be proposed for the genesis of barite

Fig. 11 Pb isotopic compositions of Sekarna galena plotted on208Pb/204Pb vs. 206Pb/204Pb and 207Pb/204Pb vs. 206Pb/204Pb diagrams(after Zartman and Doe 1981). Pb isotopic compositions of galenafrom Fedj El Adoum, Bougrine and Boujabeur deposits are fromSkaggs (2010)

Table 6 Pb isotopic analyses of galena from the Sekarna deposit

Sample references Mineral Host rocks 208Pb/206Pb 207Pb/206Pb 206Pb/204Pb 207Pb/204Pb 208Pb/204Pb

GC Galena Dolomite (unit C) 2.0725 0.8352 18.7695 15.6759 38.9004

GD1 Galena Dolomite (unit C) 2.0746 0.8362 18.7583 15.6862 38.9154

GD2 Galena Dolomite (unit C) 2.0819 0.8371 18.8627 15.7902 39.2700

GS Galena Phosphorite (unit A) 2.0729 0.8357 18.7536 15.6733 38.8752

GPH Galena Phosphorite (unit A) 2.0713 0.8348 18.7876 15.6838 38.9154

GD3 Galena Limestone (Campanian–Maastrichtian) 2.0701 0.8350 18.7418 15.6488 38.7976

GD4 Galena Limestone (Campanian–Maastrichtian) 2.0763 0.8369 18.7663 15.7046 38.9642

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in the Lower Metlaoui Formation: (1) hydrothermal input ofBa-enriched fluids and (2) early diagenetic formation withinthe sediments. The size and morphology of the barite crys-tals and their sulfur isotopic compositions may help discrim-inate between the two mechanisms.

Sekarna barite is characterized by fibroradiated concre-tion shapes in silicified phosphatic marls and is enriched byabout 4‰ relative to Eocene seawater sulfate, reflectingbarite formation from residual sulfate after preferential lossof 32S to H2S during bacterial sulfate reduction (BSR).Similar textures and isotopic patterns have been observedin Paleogene concretions in the West Carpathian Flysh inPoland (Leśniak et al. 1999), which were interpreted toreflect diagenetic processes occurring in the sediment col-umn, including BSR. Another example was presented byGoldberg et al. (2006) who described Cambrian barite–pyrite concretions in black shales of the early CambrianNiutitang Formation in China. Based on δ34S and δ18O,the authors proposed a diagenetic origin for the barite afterextensive BSR. In both studies, the enrichments in heavyisotopes were attributed to closed or semi-closed conditionsduring diagenesis in which pore water sulfate underwentextensive BSR. Barite δ34S values in the Sekarna depositsvary from the coeval seawater sulfate value to higher values.The isotopic compositions of celestite and barite are thusconsistent with data from the literature that show enrichmentby 5–10‰ relative to the associated gypsum, which isinconsistent with the simple dissolution of preexisting sul-fate minerals and reprecipitation of the sulfate as barite.

Diagenetic barite can form at the oxic–anoxic boundaryin marine sediments where BSR is also taking place. At thislocation, sulfate-reducing bacteria metabolize isotopicallylight sulfate more rapidly than isotopically heavy sulfate toproduce sulfide, leaving residual sulfate enriched in heavysulfur (Seal II et al. 2000). Thus, the enrichment in theheavy isotope of sulfur may reflect the formation of baritein restricted basins with limited water exchange with theopen ocean. In contrast, barites that form where Ba-richsolutions emanate directly into seawater should carry theisotopic signature of the Eocene seawater sulfate, which wasaround 21‰ (Claypool et al. 1980).

Sulfur source(s)

The reduced sulfur in sediment-hosted Zn–Pb ores can bederived from a variety of sources including organicallybound sulfur, H2S-bearing natural gas, contemporaneousseawater sulfate, diagenetic pyrite, and sulfate evaporitefacies in the stratigraphic section. Reduction of sulfate canbe accomplished by two processes: low-temperature bacte-rially mediated reduction (BSR) and abiotic thermochemicalreduction (TSR). BSR is common in diagenetic settings,with optimal temperatures ranging from 0°C to 70±10°C;

rates of reduction are drastically reduced at higher temper-atures (Machel 2001). TSR requires higher temperatures,>100°C, and is more efficient above 125°C (Ohmoto 1992).

At Sekarna, δ34S values for sphalerite and galena arenegative, falling in the relatively narrow range (from−16.0‰ to −9.3‰), with sphalerite δ34S values higher, onaverage, than galena values. The uniformity of the isotopiccompositions, and the fact that sphalerite δ34S is generallygreater than galena δ34S, suggests that H2S had been homog-enized in the ore-forming fluid and that sphalerite and galenaprecipitation involved fluid–mineral isotopic exchange thatapproached, although may not have attained, isotopic equilib-rium. The low (negative) δ34S values suggest that sulfatereduction occurred in locations, perhaps pore waters of anoxicsediments, that were in communication with a large sulfatereservoir perhaps overlying seawater.

The high δ34S value of Sekarna barite (25.1±0.2‰) isgenerally consistent with a seawater origin for the sulfate(Claypool et al. 1980). The δ34S of Palaeogene seawatersulfate, which has been determined from the analyses ofevaporates and marine barite, was about 21‰ (Claypool etal. 1980; Paytan et al. 1998). Gypsum precipitated fromseawater will have an isotopic composition that is higherthan the dissolved sulfate by 1.65±0.12‰ (Thode andMonster 1965). There is no significant fractionation associ-ated with the dissolution of sulfates, but repeated dissolutionand reprecipitation of older sulfates could lead to a signifi-cant increase in δ34S relative to the marine value. We preferan interpretation in which the high Sekarna barite values(average025.1‰, n02) reflect isotopic enrichment of sul-fate as a consequence of BSR.

The isotopic fractionation associated with sulfate reduc-tion that produced the Sekarna sulfides was about 33‰(21‰ difference between Eocene marine sulfate (Ayora etal. 1995; Paytan et al. 1998) and ore sulfides (average0−11.8‰)). This fractionation is close to the average differ-ence between seawater sulfate and sedimentary sulfides ofthe same age (40‰; Ohmoto 1986). The similarity of bariteδ34S values to the Tertiary marine sulfate value suggests thatbarite sulfur was derived from Tertiary evaporite minerals orsulfate contained in pore water in surrounding sedimentaryunits.

Phosphorite-hosted mineralization displays textural evi-dence that sulfides replaced earlier barite (barite is found assubmicroscopic inclusions in sphalerite and galena). Thus,sulfur for ore formation may have been derived from areductive dissolution of barite as well as from Eocene evap-orites or seawater. Diagenetic pyrite was not a major sulfursource. Although anoxic sediments of the Cenomanian–Turonian Bahloul Formation are strongly enriched in diage-netic pyrite, this sulfur is much lighter, with δ34S values of−18±2‰ (Peevler et al. 2003). More deeply buried Triassicevaporites could also have been a source of sulfur, as

Miner Deposita (2012) 47:545–562 559

proposed for several deposits in northern Tunisia (Sheppardand Charef 1990; Orgeval 1994; Sheppard et al. 1996;Bouhlel 1993, 2005, Bouhlel et al. 2007). Reduction ofdeeply buried marine sulfate is commonly interpreted toinvolve TSR and to result in positive δ34S values for oreminerals. However, it is important to note that positiveisotopic values can also result from closed-system BSR, aswell as from the mixing of sulfur from multiple sourcesinvolving multiple reduction processes.

Fluid inclusion temperatures within the ore zone arepermissive of TSR acting on local sulfate, but the low anduniform sulfide δ34S values at Sekarna suggest a homoge-neous fluid source and do not support the involvement ofTSR in ore genesis. It is unlikely that sphalerites containingboth the moderate-salinity and high-salinity inclusion typesprecipitated from two distinct fluids. Stretching during post-entrapment overheating can be excluded because the low-salinity fluid inclusions show the highest homogenizationtemperatures. The temperature of sulfide depositionexceeded the temperature range at which bacterial processesare efficient. BSR must have occurred at a different placeand/or time relative to sulfide deposition. It is quite possiblethat sulfate was reduced by bacterial processes, and then theresulting H2S mobilized by higher temperature brines toform sphalerite and galena. Reaction between dissolvedH2S and metallic chloride complexes in mixing and diffu-sion zones resulted in rapid sulfide precipitation.

Conclusions

The Sekarna sulfide deposit occurs within an organic-richphosphatic host rock that differs from the host rocks formost MVT deposits. The deposit shows predominantly latediagenetic to epigenetic styles of mineralization thatreplaced carbonate cement. Mineral assemblages are simpleand are dominated by sphalerite, galena, Fe sulfides, andminor barite. The ore-forming fluids were brines of moder-ate salinity and temperature. The minor element geochem-istry of Sekarna galena and sphalerite resembles the minorelement geochemistry of sulfides in carbonate-hosted MVTdeposits elsewhere in the world.

The sulfur isotopic compositions of Sekarna sulfides arerelatively uniform; the data are interpreted to reflect a singlesulfur source that most likely involved BSR acting on Eo-cene marine sulfate. The Pb isotopic compositions of galenaare also relatively uniform. The data suggest that the orefluids circulated through deep levels of the upper crust.

Similarities to MVT deposits include the simple miner-alogy, late diagenetic to epigenetic timing of mineralization,sulfide replacement of carbonate, and the involvement ofbasinal brines. However, the uniform sulfur and Pb isotopic

compositions observed at Sekarna, which imply a singlemetal source, are atypical of MVT deposits.

Acknowledgments This paper is part of a Doctorate thesis of thefirst author. We would like to express our thanks to Sheldon A. Skaggs(Department of Geology, Building University of Georgia Athens) forlead isotope analyses. We wish to thank Karen Duttweiler Kelley(USGS, Denver USA) for constructive comments and for English edits.We are grateful to Christian Marignac for the editorial handling and hiseffort to improve this article. Yves Fuchs and Etienne Delloul, tworeviewers of this article, are gratefully acknowledged for their criticaland constructive comments that helped significantly improve the man-uscript. Bernd Lehmann, editor-in-chief of the journal, is thanked forhis constructive comments on the article. S. Bouhlel gratefullyacknowledges support by Fulbright Scholar Grant that made possiblethe carbon, oxygen, and sulfur analysis at the USGS Denver, USA.

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